The Serinhisar-Acıpayam basin of western Anatolia hosts a Neogene alkaline lake which formed in some graben and semi-graben depression zones as a result of N–S tension. The basin is filled with fluvial and lacustrine sediments dominated by clayey materials. The filling of the basin with fine sediments and associated water level changes caused the development of swampy and/or semi-swampy, alkaline-lake environments where sepiolite, palygorskite, saponite and dolomitic sepiolite or palygorskite precipitated periodically in the basin. Sepiolite is predominant in the Kuyucak section and is intercalated with saponite-dominated levels, whereas saponite accompanied palygorskite at Kocapınar where basaltic volcanism occurred. The contacts between sepiolite-palygorskite and saponite levels are more or less sharp, reflecting rapid changes in the physicochemical conditions of the depositional environment. Micromorphological images reveal that both sepiolite and palygorskite grew as interwoven fibers or fiber bundles and masses where dolomite was absent, indicating direct precipitation from solution, whereas fibrous networks grew authigenically on and out of dolomite in dolomitic sepiolite and dolomitic palygorskite. Saponite is either green or reddish brown due to its organic material-rich content and derivation from products of basaltic volcanism. Synsedimentary basaltic volcanism was the main source of Fe and Al, whereas Si and Mg were derived from surrounding ultrabasic and detrital units and partly from the volcanism. It can be concluded that sepiolite, palygorskite and saponite formed either by direct precipitation from alkaline lake water or authigenically from interstitial pore-water between dolomite rhombs as controlled by concentration of Si, Mg, Al and Fe, rather than by mutual transformation.

Due to their unique polarity, pore-size distribution, and high surface areas, pillared and delaminated clays are potentially useful materials for the adsorption of environmental toxicants. To determine their properties for adsorption of chlorinated phenols, alumina-pillared montmorillonite (APM), chromia-pillared montmorillonite (CPM), and alumina-delaminated Laponite (ADL) were reacted with aqueous pentachlorophenol (PCP) solutions in batch equilibrium experiments. An hydroxy-Al Laponite (HAL) in which the Na+ exchange ions were replaced by ions of the type Al13O4(OH)(24+x)(H2O)(12-x)(7-x)+ was included in the study. With ADL as the adsorbent, the extent of PCP adsorption increased with decreasing pH, and then became constant at pH ≤ pKa. Thus, the neutral phenol was preferred over the phenolate form. Binding of neutral PCP at pH 4.7 to all adsorbents never reached saturation values, and the loadings achieved were limited by the water solubility of the adsorbate. Among the pillared and delaminated clays investigated, ADL exhibited the largest capacity for physical adsorption of PCP at pH 4.7. Differences in the PCP binding capacities for APM, CPM, and ADL suggested that adsorption was dependent on the pore structure and surface composition of the modified clay adsorbent, not on surface area alone. HAL exhibited quantitative uptake of PCP at the 8 µmole/g level, indicating that a chemi-sorption mechanism may operate for PCP binding to this adsorbent. Adsorption of 3-chlorophenol, 3,5-dichlorophenol, and 3,4,5-trichlorophenol by ADL at pH 7.4 increased as the degree of hydrophobicity and chlorination of the phenol increased; hence, the binding capacity was not limited by the molecular size of the adsorbate. In contrast to the adsorption properties observed for pillared, delaminated, and hydroxy-interlayered clays, Na+-montmorillonite and Na+-Laponite did not adsorb PCP from aqueous solution.

Mössbauer spectra of 15 smectites were investigated. In these samples, ferric iron occupies both M(1) and M(2) octahedral sites, the distribution being partly determined by the relative covalency of the bonds formed.

The quadrupole splittings are linearly related to b−3. They show that Fe3+ octahedra are much more distorted in montmorillonite and beidellite than in nontronite and volkonskoite and that M(1) sites are more prone to change than M(2). Ferrous iron occurs in relatively undistorted octahedra in some otherwise distorted octahedral sheets and vice versa.

A process for treating smectite-bearing rock samples that utilizes LR White resin; allows TEM observation of expanded smectite interlayers and therefore preservation of original rock textures. Examples of several lattice fringe images are shown, including: (1) Dioctahedral smectite layers from a shale (1388.9 meter depth, Texas Gulf Coast) give fringes that consistently have spacings of 1.2–1.3 nm, yet duplicate other features previously observed in collapsed samples. (2) Packets of illite layers give fringes with 1.0-nm spacings coexisting with packets of (dominantly) R1 I/S having 2.1-nm lattice fringe spacings in a Gulf Coast shale from 4742.1 m. (3) Rectorite from Garland Co., Arkansas gives 2.3-nm lattice fringes. Samples with wide ranges of I/S ratios and lithologies have been found to be permanently expanded with retention of original textures, commonly leading to unambiguous identification of illite and smectite interlayers in lattice fringe images.

Polytypism in trioctahedral 1 : 1 phyllosilicates results from two variable features in the structure. (1) The octabedral cations may occupy the same set of three positions throughout or may alternate regularly between two different sets of positions in successive layers. (2) Hydrogen bonding between adjacent oxygen and hydroxyl surfaces of successive layers can be obtained by three different relative positions of layers: (a) direct superposition of layers, (b) shift of the second layer by a/3 along any of the three hexagonal X-axes of the initial layer, with a positive or negative sense of shift determined uniquely by the octahedral cation set occupied in the lower layer, and (c) shift of the second layer by ± b/3 along Y1 (normal to X1) of the initial layer regardless of octahedral cation sets occupied. Assuming ideal hexagonal geometry, no cation ordering, and no intermixing in the same crystal of the three possible types of layer superpositions, then twelve standard polytypes (plus four enantiomorphs) with periodicities between one and six layers may be derived. Relative shifts along the three X-axes lead to the same layer sequences derived for the micas, namely 1M, 2M1, 3T, 2M2, 2Or, and 6H. Polytypes 1T and 2H1 result from direct superposition of layers. Layer shifts of b/3 lead to polytypes designated 2T, 3R, 2H2, and 6R. The twelve standard 1 : 1 structures can be divided into four groups (A = 1M, 2M1, 3T; B = 2M2, 2Or, 6H; C = 1T, 2T, 3R; D = 2H1, 2H2, 6R) for identification purposes. The strong X-ray reflections serve to identify each group and the weaker reflections differentiate the three structures within each group. Examples of all four groups and of 9 of the 12 individual structures have been identified in natural specimens. Consideration of the relative amounts of attraction and repulsion between the ions in the structures leads to the predicted stability sequence group C > group D > group A > group B, in moderately good agreement with observed abundances of these structural groups.

Imogolite is a tubular aluminosilicate which is common in Andosols and Spodosols. The high pH at point-of-zero charge at the outer parts of the tube and the anomalously high chloride adsorption of imogolite suggested that there may be structural charge associated with this mineral. The structural charge may arise because of changes in bond valence imposed by the incorporation of orthosilicate anions in a gibbsite-type sheet. By using a Basic Stem Model approach, it is shown that the surface charge properties of imogolite are explained if the mean Al-O bond valence of the outer -Al2OH groups is higher than the inner -Al2OHSiO3 groups. Hence, a weak positive charge is developed on the outer tube walls whereas a negative charge develops in the tubular pores. The best model fits were obtained where either one or two units of structural charge per unit cell of tube were assumed. The model may also explain why imogolite tubes are normally aggregated in large bundles in close hexagonal packing, because bound counterions may hold the tubes together. However, to arrive at good model descriptions, the deprotonation of -Al2OH groups must occur at a higher pH than that expected when assuming that all surface oxygens form two hydrogen bridges with H2O. A more precise structure of imogolite is required to test fully this hypothesis.

The hydroxyl deformation modes of kaolins have been studied by Fourier transform (FT) Raman spectroscopy. Kaolinites showed well-resolved bands at 959, 938 and 915 cm−1 and an additional band at 923 cm−1. For dickites, well-resolved bands were observed at 955, 936.5, 915 and 903 cm−1. Halloysites showed less-resolved Raman bands at 950, 938, 923, 913 and 895 cm−1. The first 3 bands were assigned to the librational modes of the 3 inner-surface hydroxyl groups, and the 915-cm−1 band was assigned to the libration of the inner hydroxyl group. The band in the 905 to 895 cm−1 range was attributed to “free― or non-hydrogen-bonded inner-surface hydroxyl groups. The 915-cm−1 band contributed ~65% of the total spectral profile and was a sharp band with a bandwidth of 11.8 cm−1 for dickite, 14.0 cm−1 for kaolinites and 17.6 cm−1 for halloysites. Such small bandwidths suggest that the rotation of the inner hydroxyl group is severely restricted. For the inner-surface hydroxyl groups, it is proposed that the hydroxyl deformation modes are not coupled and that the 3 inner-surface deformation modes are attributable to the three OH2-4 hydroxyls of the kaolinite structure. For intercalates of kaolinite and halloysite with urea, a new intense band at ~903 cm−1 was observed with concomitant loss in intensity of the bands at 959, 938 and 923 cm−1 bands. This band was assigned to the non-hydrogen-bonded hydroxyl libration of the kaolinite-urea intercalate. Infrared reflectance (IR) spectroscopy confirms these band assignments.

Fougerite (IMA 2003-057) is a mixed M(II)-M(III) hydroxysalt of the green rust group, where M(II) can be Fe or Mg, and M(III) is Fe. The general structural formula is: ${[{\text{Fe}}_{1-}^{2+}x{\text{Fe}}_x^{3+}{\text{Mg}}_y{(\text{OH})}_{2+2y}]}^{+x}{[x\text{/}n{A}^{-n}.m{\text{H}}_2\text{O}]}^{-x}$ where A is the interlayer anion and n its valency, with 1/4 ≼ x/(1+y) ≼ 1/3 and m ≼ (1−x+y). The structure of green rusts and parent minerals can accommodate a variety of anions, such as OH−, Cl−, ${\text{CO}}_3^{2-},{\text{SO}}_4^{2-}$${\rm{CO}}_3^{2 - },\;{\rm{SO}}_4^{2 - }$. The structure of the mineral was studied by Mössbauer, Raman and X-ray absorption spectroscopies (XAS) at the FeK edge. Mössbauer spectra of the mineral obtained at 78 K are best fitted with four doublets: D1 and D2 due to Fe2+ (isomer shift δ ≈ 1.27 and 1.25 mm s−1, quadrupole splitting ΔEQ ≈ 2.86 and 2.48 mm s−1, respectively) and D3 and D4 due to Fe3+ (δ ≈ 0.46 mm s−1, ΔEQ ≈ 0.48 and 0.97 mm s−1, respectively). Microprobe Raman spectra obtained with a laser at 514.53 nm show the characteristic bands of synthetic green rusts at 427 and 518 cm−1. X-ray absorption spectroscopy shows that Mg is present in the mineral in addition to Fe, that the space group is and the lattice parameter a ≈ 0.30–0.32 nm. The mineral forms by partial oxidation and hydrolysis of aqueous Fe2+, to give small crystals (400–500 nm) in the form of hexagonal plates. The mineral is unstable in air and transforms to lepidocrocite or goethite. The name is for the locality of the occurrence, a forested Gleysol near Fougères, Brittany, France. Its characteristic blue-green color (5BG6/1 in the Munsell system) has long been used as a universal criterion in soil classification to identify Gleysols. From a thermodynamic model of soil-solution equilibria, it was proposed that for the eponymous mineral, Fougères-fougerite, OH− may be the interlayer anion. In other environments, the interlayer anion may be different, and other varieties of fougerite may exist. Fougerite plays a key role in the pathways of formation of Fe oxides.

The southern San Joaquin Valley contains more than 7 km of sedimentary fill, largely Miocene and younger in age. Ancient depositional environments ranged from alluvial fans at the basin margins to turbidite fans toward the basin center. Mixed-layer illite/smectite (I/S) dominates the <2-μm fraction of Miocene shales, and kaolinite is abundant in Miocene sandstones. I/S from carbonate-cemented sandstones contains 5–20% more smectite layers than I/S from uncemented sandstones. The timing of cementation correlates with the proportion of smectite layers in the I/S, suggesting that cementation slowed the illitization process. Smectite and I/S with > 80% expandable layers occur at present burial temperatures of 120°-l 40°C in Miocene sandstones and shales. This highly expandable I/S is restricted to areas covered by thick deposits (1000-2500 m) of Pleistocene sediments. Rocks of similar age and at equivalent temperatures, but covered by <900 m of Pleistocene sediments, contain I/S having low expandabilities (<30%).

Microprobe analyses of 16 discrete smectite and smectite-rich I/S clays indicate an average montmorillonite composition of:

$$\left( {C{a_{0.2}}{K_{0.15}}N{a_{0.1}}} \right)\left( {A{l_{3.0}}F{e^{3 + }}_{0.35}M{g_{0.75}}T{i_{0.03}}} \right)\left( {S{i_{7.7}}A{l_{0.3}}} \right){O_{20}}{\left( {OH} \right)_4} \cdot n{H_2}O.$$

Residence time at different temperatures appears to be an important influence on the percentage of smectite layers in I/S from the San Joaquin basin. Areas containing I/S with high expandabilities (e.g., 95% smectite layers) have a time-temperature index (TTI) of 4.0-4.5 at 120°C, whereas areas containing I/S with low expandabilities (e.g., 30% smectite layers) have a TTI of 5.0. Present data suggest that highly expandable I/S changed to slightly expandable I/S over a narrow temperature interval (10°-20°C). Differences in the potassium availability from detrital components and in the K+/H+ activity ratios of pore water do not appear to be related to the differences in the percentage of smectite layers of these I/S clays.

The cation exchange capacity (CEC) is one of the most important properties of clays in terms of their performance in both natural and technical processes. For decades, common methods for determining exchangeable cations have failed when calcareous clays or soils were examined, because calcite is at least partly dissolved throughout the exchange experiment which in turn increases measureable Ca2+ concentrations. As a result, exchangeable Ca2+ and the sum of exchangeable cations by far exceed the amount of negative charges. In the past, the silver-thiourea method (AgTU) has been modified to overcome this problem (AgTUcalcite), but is unsatisfactory as the method is laborious. In the present study three new methods based on two alternative metal complexes, cobalt(III) hexamine (CoHex) and copper(II) triethylenetetramine (Cu-trien), are proposed. The optimum solid/liquid ratios of these methods and the optimum complex concentration of Cu-trien are reported, depending on the mineralogical composition of the samples. The key development is that the exchange solutions are saturated with respect to calcite prior to the experiment. Approximately 70–90% of the dissolution of calcite present as an admixture in a clay sample is suppressed in the subsequent cation exchange experiment, but not all. The Ca2+ exchange is not suppressed and there is no evidence for any precipitation of this Ca2+. Three possibilities for how to handle this problem are discussed, one of which is to perform no further correction. The resulting error arises from the remaining calcite solubility of the different solutions after pre-treatment with calcite. This corresponds to errors of 0.2–1.3 (CoHexcalcite)and 0.7–8.4 (Cu-triencalcite) meq/100 gCa2+ for samples with small and large CEC values, respectively. As a consequence of the poor performance of Cu-triencalcite for samples with large CEC, a more concentrated Cu-trien5 × calcite solution was developed which performed much better: 0.1–0.8 meq/100 g(Cu-trien5 × calcite). For Cu-trien5 × calcite and CoHexcalcite at least, the errors are in the range of the non-systematic scattering for exchangeable Ca2+ determination. Therefore, the methods suggested provide ‘operationally correct’ Ca2+ values without additional effort. Moreover, owing to the high selectivity of the index cations applied in the present study, only one exchange step is required, providing a significant advantage over the AgTUcalcite method.

Berthierine (formerly chamosite) occurs as concretions, lenses, and bands in carbonaceous, kaolinitic shale of freshwater coal-swamp deposits in Paleogene and Upper Triassic coal measures of Japan. Textural relations in thin sections of the Triassic berthierine rocks and a siderite-kaolinite-berthierine-quartz assemblage in Paleogene rocks indicate that the berthierine formed by reaction of siderite with kaolinite. The transformation of siderite and kaolinite to berthierine and quartz occurs progressively under reducing conditions between 65° and 150°C and at burial depths of 2–5 km. Utatsu berthierine is an aluminous, low-Mg variety as compared with berthierine pellets in modern marine and estuarine sediments and in ancient marine ironstones. Fe is the dominant octahedral cation with Fe2+ ≫ Fe3+. The composition of the berthierine varies between different morphological types. Utatsu berthierine transformed to ferrous chamosite when kaolinite in the host shale changed to pyrophyllite. These transformations are estimated to have occurred at ∼160°C and at a burial depth of ∼3 km.

Sets of replicate hydraulic conductivity tests were conducted using 100 mM KCl and 20 and 40 mM CaCl2 solutions to evaluate how changes in hydraulic conductivity are related to changes in the exchange complex and physical properties (water content and free swell) of prehydrated and non-prehydrated bentonite used for geosynthetic clay liners (GCLs). Essentially no change in hydraulic conductivity and water content (or void ratio) occurred during tests with the 100 mM KCl solution even though K+ was replacing Na+ on the exchange complex. In contrast, for the CaCl2 solutions (20 mM and 40 mM), the hydraulic conductivity increased and the free swell and water content decreased as exchange of Ca2+ for Na+ occurred. Faster changes in hydraulic conductivity and the exchange complex occurred in the tests with the 40 mM CaCl2 solution and the non-prehydrated bentonite (i.e. the hydraulic conductivity changed more rapidly when exchange occurred more rapidly). Even though exchange of Ca2+ for Na+ was essentially complete at the end of testing, the hydraulic conductivity obtained with the 20 mM CaCl2 solution was less and the water content greater than obtained with the 40 mM CaCl2 solution (2.6 × 10−8 cm/s vs. 6.7 × 10−8 cm/s, 122% vs. 111%, and 3.2 vs. 2.9). Similarly, the prehydrated bentonite had lower hydraulic conductivity (1.9 × 10−8 cm/s vs. 6.7 × 10−8 cm/s) and greater water content (167% vs. 111%) than the non-prehydrated bentonite at the end of testing, even though Ca-for-Na exchange was essentially complete.

In an attempt to resolve the structure of opal-CT and opal-C more precisely, 24 opal samples from bentonites, Fuller's Earths, zeolite tuffs, biogenic silicas and silicified kaolins have been analyzed by high resolution transmission electron microscopy (HRTEM) and X-ray diffraction (XRD). Results of this examination demonstrate that opal-C and opal-CT are part of a continuous series of intergrowths between end-member cristobalite and tridymite stacking sequences.

These findings are consistent with Flörke's (1955) interpretation of the most intense opal peak at ~4 Å as a combination of the (101) cristobalite and (404) tridymite peaks. The position and width of this peak are controlled by the relative volume of the two stacking types and the mean crystallite size. Direct evidence obtained by HRTEM provides data showing various stacking sequences in opals. Broadening due to crystallite size alone was determined by directly measuring crystallite size by TEM and comparing the measured size to the apparent size calculated using the Scherrer equation. XRD peak broadening is also described in terms of various contributions from structural disorder. The mean opal crystallite size ranges from 120 to 320 Å. For samples at either end of the size range, the crystallite size plays a larger role, relative to stacking disorder, in controlling peak broadening.

The dehydration and rehydration processes of the clay mineral palygorskite (PFl-1) were studied by textural analysis, thermogravimetric analysis connected with mass spectrometry (TGA-MS), and 29Si and 1H solid-state NMR techniques. The TGA-MS results clearly reveal weight losses at maxima of 70°C, 190°C, 430°C and 860°C. PFl-1 is characterized by a micropore area of 93 m2/g, corresponding to a micropore volume of 47 mm3/g. These values are also obtained for the sample heated up to 200°C for 20 h. Further heating at 300°C produces a collapse of the structure, as shown by the almost complete loss of microporosity.

The 29Si NMR spectra of palygorskite show two main resonances at −92.0 and −97.5 ppm, attributed to one of the two pairs of equivalent Si nuclei in the basal plane. A minor resonance at −84.3 ppm is attributed to Q2(Si-OH) Si nuclei. The resonance at −92.0 ppm is assigned to the central Si position, while the resonance at −97.5 ppm is assigned to the edge Si sites. It is confirmed bysolid-state 29Si and 1H NMR that nearly complete rehydration is achieved by exposing palygorskite samples that have been partially dehydrated at 150°C and 300°C, to D2O or water vapor at room temperature. When the rehydration is accomplished with D2O, the atoms are disordered across all the protons sites.

The conversion of smectite to illite by way of a mixed-layer illite/smectite (I/S) series was found to be the major depth-related reaction in clay-mineral assemblages from two cored sedimentary sequences in the Rocky Mountains. The I/S reaction occurred in both interbedded sandstone and shale of Upper Cretaceous and lower Tertiary age in the Green River basin, Wyoming, and in chalk and chalky shale of the Upper Cretaceous Niobrara Formation, Denver basin, Colorado. As the proportion of illite layers in I/S increased with depth in these rocks, the amount of I/S in the clay fraction decreased, and the amount of discrete illite increased. Scanning electron microscopy revealed that the morphologies of highly expansible, randomly interstratified I/S clay (samples from shallow cores) exhibited no distinctive intergrowth or overgrowth textures. In deeply buried rocks containing highly illitic, ordered I/S and abundant discrete illite, however, fibers or laths of illite were formed on earlier I/S substrates. Less commonly, I/S of low expandability shows morphological features of both smectite and illite whereby rigid laths of illite appear to have formed diagenetically from the wall surfaces of I/S honeycombs. This combined morphology suggests some dissolution and reprecipitation (or some reorganization) of materials from the I/S substrate as the substrate was transformed into a more illitic mixed-layer clay.

These data suggest that some I/S clay was destroyed by the selective cannibalization of smectite layers in I/S to provide the components needed to make a more illitic I/S. Moreover, discrete illite and other minerals apparently were formed during the reaction. Also, coarser mineral phases, such as potassium feldspar and detrital micas, may not have been required to supply the chemical components in the reaction. The observations provide an explanation for late diagenetic I/S reactions that occurred in restricted or relatively closed geochemical systems, as in early cemented rocks having extremely low permeability and little or no potassium feldspar.

Fourier transform infrared (FTIR) spectroscopy investigations in the near infrared (N1R) region of synthetic and natural kaolinites with various octahedral substitutions have been carried out in order to elucidate the relationships between the substituted cations and specific features of the NIR spectra. The combination modes of the OH stretching and bending vibrations characterizing Fe(III), Ga(III) and Cr(III) octahedral substitutions are identified in the NIR region at 4466, 4498 and 4474 cm-1, respectively, and the first overtones of the OH stretching vibrations at 7018, 7018 and 6986 cm-1, respectively. As far as we know, the bands of kaolinites containing Ga(III) or Cr(III) have not been reported yet. For both Ga(III) and Cr(III), the NIR observations explain why the bending vibration bands of AlGaOH and AlCrOH groups are not observed in the middle infrared (MIR) region.

Mössbauer spectroscopy of dioctahedral phyllosilicates showed that on dehydroxylation iron which originally occupied M(2) and M(l) sites became, respectively, 5- and 6-coordinated. The 6-coordinated sites are very distorted. No migration of cations occurs in the course of heating the specimens for 1–3 hr at 600°–700°C.

By using a combination of several physicochemical methods, different successive stages of the dehydroxylation process could be distinguished: (1) migration of protons; (2) localized dehydroxylation of individual associations without significant change in the overall configuration of the octahedral sheets; and (3) loss of most of the hydroxyl groups with concomitant changes in the cell dimensions. Penetration of Li into the octahedral sheets does not affect the course of the reaction, but reduces the dehydroxylation temperature and the stability of the products.

Dehydroxylation was preceded by or associated with the oxidation of any divalent iron present. Fe3+ derived from Fe2+ was indistinguishable by Mössbauer spectroscopy from iron initially present in the trivalent form. High concentrations of Fe lower the dehydroxylation temperature and reduce the stability of the dehydroxylate to the extent that partial disintegration may precede complete dehydroxylation.

Chemical activity diagrarns, prepared to illustrate the properties expected if mixed-layer mont-morillonite/illite is regarded as a solid solution, are compared to those derived from a treatment of these materials as a mixture of two phases. If the system is a solid solution, the coexisting aqueous solution should range from higher dissolved silica contents in the presence of kaolinite and a montmorillonite end member to lower dissolved silica in the presence of kaolinite and an illitic end member. Silica concentration in the aqueous solution might vary by a factor of as much as six. If the system is two phase, the silica content of a solution in equilibrium with kaolinite and both phases would be fixed at a given T and P, as would a solution equilibrated with both phases and K-feldspar. In the absence of a third phase, silica in equilibrium with both phases should be nearly constant, but increase with increasing ratio of K+/H+ in solution. Available data on coexisting aqueous solutions apparently are more nearly consistent with two phases than with a solid solution.